Earth and Planetary Science Letters 286 (2009) 110–121
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Earth and Planetary Science Letters
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l
Late Paleocene high Laramide ranges in northeast Wyoming: Oxygen isotope study of
ancient river water
Majie Fan ⁎, David L. Dettman
Department of Geosciences, University of Arizona, Tucson, Arizona 85721, United States
a r t i c l e
i n f o
Article history:
Received 26 September 2008
Received in revised form 16 June 2009
Accepted 16 June 2009
Available online 19 July 2009
Editor: T.M. Harrison
Keywords:
Rocky Mountains
paleoelevation
oxygen isotope ratios
Laramide
freshwater bivalve
a b s t r a c t
The distribution and initial timing of the establishment of high surface elevations in the Rocky Mountains
during the Early Cenozoic remain controversial despite the importance of these data in testing tectonic models
for this region. We track the timing and distribution of high elevation in the Rocky Mountains during the Late
Cretaceous–Early Eocene by examining annual and seasonal δ18O values of the ancient river water, which are
calculated from the δ18O values of well-preserved freshwater bivalve shells. In the Powder River basin of the
eastern Laramide province, the δ18O values of the ancient river water vary between − 23.0‰ and − 8.0‰SMOW
in both seasonal and annual records in the Late Paleocene–Early Eocene. The large variation suggests that the
ancient rivers were fed yearly or seasonally by snowmelt from highlands of 4.5 ± 1.3 km. This can be explained
by the existence of the Bighorn Mountains and Black Hills with a drainage pattern similar to the present in
northeast Wyoming. The δ18O values of ancient river water along the front of the Sevier thrust belt generally
follow a trend from lower values in north, −14.2 ± 1.4‰ in the Early Paleocene Crazy Mountains basin, to
higher values in south, − 11.1 ± 0.8‰ in the Late Paleocene Bighorn basin, and − 7.1 ± 1.6‰ in the Early Eocene
Washakie basin. The variations within each basin are relatively small. These rivers most likely rise in the Sevier
thrust belt, and may reflect highland elevation of 1–2 km. The δ18O values in the Alberta foreland and Williston
basin are very low (− 20.5‰) in the Late Cretaceous, indicating the rivers were fed by snowmelt from the
Canadian Rocky Mountains of 4.3 ± 1.0 km high. The attainment of high elevation in the eastern Laramide
province prior to the western province could be explained by southwestward progression of back-thrusts soled
into an earlier east-directed master detachment, which may be formed by the westward rollback of subducted
shallow slab.
Published by Elsevier B.V.
1. Introduction
The Laramide orogeny of the Rocky Mountains is a system of
basement-cored uplifts and intervening basins that formed ~80–
40 Ma in the foreland basin of the Sevier thrust belt in the western
United States (Dickinson and Snyder, 1978; Bird, 1998; DeCelles, 2004).
Analogous to modern flat-slab subduction in western South America
(Jordan and Allmendinger, 1986), Laramide uplifts are the result of the
NE–SW compression due to shallow subduction of the Farallon Plate
beneath the North American Plate (e.g., Dickinson and Snyder, 1978;
Bird, 1998; Saleeby, 2003; DeCelles, 2004). Although this region is very
well studied, fundamental questions remain unanswered: how does
shallow subduction thicken foreland crust and produce a landscape
with intervening basins and ranges, and how is basement-involved
deformation connected to the thin-skinned Sevier fold and thrust belt.
A range of tectonic models have been proposed to explain the
mechanisms of Laramide deformation: basal traction (Bird, 1998),
lithospheric buckling (Tikoff and Maxson, 2001), intracrustal flow
⁎ Corresponding author. Tel.: +1 520 621 6014; fax: +1 520 621 2672.
E-mail address: [email protected] (M. Fan).
0012-821X/$ – see front matter. Published by Elsevier B.V.
doi:10.1016/j.epsl.2009.06.024
(McQuarrie and Chase, 2001), and thrust, back-thrust and crustal
detachment (Erslev, 1993, 2005). The timing, magnitude, and pattern
of high surface topography in the Laramide province need to be
examined to help answer these basic questions and to test tectonic
models.
Oxygen isotope ratios of precipitation derived from authigenic
minerals (e.g., paleosol carbonate, biogenic apatite and aragonite)
have been applied as paleoaltimeters in numerous studies (e.g.,
Dettman and Lohmann, 2000; Garzione et al., 2006; DeCelles et al.,
2007; Quade et al., 2007). This approach is based on the decrease in
oxygen isotope ratios of precipitation as elevation increases, which is
controlled by progressive condensation of atmospheric water vapor
due to cooling or adiabatic expansion as the vapor mass ascends,
leading to Rayleigh isotope fractionation (Dansgaard, 1953; Rowley,
2007). At present, there have been only a few studies addressing the
paleoelevation of the Laramide Rocky Mountains (Gregory and Chase,
1992; Norris et al., 1996; Wolfe et al., 1998; Dettman and Lohmann,
2000; Fricke, 2003; Sewall and Sloan, 2006). Even though these studies
have been mostly focused on Eocene elevations, they have yielded
highly varying results. By using leaf margin analyses, Wolfe et al.
(1998), and Gregory and Chase (1992) suggested that the Eocene
M. Fan, D.L. Dettman / Earth and Planetary Science Letters 286 (2009) 110–121
111
Fig. 1. Shaded relief map of the central and Canadian Rocky Mountains. Symbols denote sampling locations and mean δ18O values of fossil shells. Basins discussed in text are bordered
by dotted line. AB: Alberta foreland basin; BH: Bighorn basin; BlH; Black Hills; BM: Bighorn Mountains; BT: Beartooth Mountains; CMB: Crazy Mountains basin; GR: Green River
basin; PD: Powder River basin; SC: Sage Creek basin; WR: Wind River basin; WRR: Wind River Range; WaB: Washakie basin; WiB: Williston basin. Arrows are paleocurrent directions
(Flores and Ethridge, 1985; Dickinson et al., 1988; DeCelles et al., 1991; Borrell and Hendrix, 2000; Mack and Cole, 2005).
Fig. 2. Generalized stratigraphic columns of sedimentary successions involved in each studied basin (age determinations of strata with sampled fossil unionid are listed in Table 1).
Key references: Cherven and Jacob, 1985; Flores and Ethridge, 1985; Lerbekmo and Coulter, 1985; Robinson and Honey, 1987; Dickinson et al., 1988; DeCelles et al., 1991; Eberth and
Hamblin, 1993; Buckley, 1994.
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Table 1
Age constraints for the study intervals in each basin.
Sample
location
Formation
Alberta foreland
basin 1
Alberta foreland
basin 2*
Williston basin
Oldman and Dinosaur Park
Formations
Horseshoe Canyon Formation
Crazy Mountain
basin
Bighorn basin
Bear, Lebo, and Melville Member of
Fort Union Formation
Fort Union and Willwood Formations
Powder River
basin
Washakie basin
Tongue River Member of Fort Union
and Wasatch Formations
Luman Tongue Member of Green
River Formation
Hell Creek Formation and Tullock
Member of Fort Union Formation
Age (Ma) Constraint
74–76
68–73
62.3–68
61–65
53.6–59
54–58
53–54
Dating method
Bentonite bed above the Dinosaur Park–Oldman K–Ar, Ar–Ar
Formation contact
C31r.1r–C32r.1r
Magnetostratigraphy
K–T boundary, bentotite beds
Interval of C26r–C29r,
Reference
Eberth and Hamblin (1993)
Lerbekmo and Coulter (1985)
(Swisher et al., 1993; Lund et al.,
Ar–Ar,
Magnetostratigraphy, 2002)
NALMA
Magnetostratigraphy (Buckley, 1994; Butler et al., 1987)
Interval of C24n–C26r, carbon isotope excursion Ar–Ar and
Magnetostratigraphy
Coal stratigraphy, occurrence of Platycarya
NALMA,
pollen , and correlation with Bighorn Basin
Biostratigraphy
Volcanic tuff above the Member, top of C24n
Ar–Ar and
Magnetostratigraphy
(Koch et al., 1995; Secord et al.,
2006)
(Wolberg, 1979; Robinson and
Honey, 1987; Wing et al., 1991)
Smith et al. (2003)
Note: *Only one shell bed in the locality.
southern Rocky Mountains were more than 3 km high, similar to
today's elevation. This agreed with the δ18O values of lake microbial
carbonates in the Green River basin (Norris et al., 1996), the δ18O
values of unaltered freshwater bivalves (Dettman and Lohmann,
2000), and regional GCM modeling of Early Paleogene Laramide
foreland (Sewall and Sloan, 2006). However, Morrill and Koch (2002)
concluded that diagenesis may have altered the δ18O of lacustrine
microbial carbonates in the Green River basin. Moreover, Fricke (2003)
argued that Early Eocene Laramide range elevations were ~500 m
based on similar δ18O values of mammal teeth from three Wyoming
basins. These conflicting conclusions may arise from a number of
factors: 1). diagenetic overprinting of original isotopic patterns; 2). not
accounting for all the non-altitude factors that can affect the δ18O of
surface water (e.g., latitude, temperature); 3). small sample numbers
failing to document a regional pattern in surface water δ18O values.
In this study we survey the oxygen isotope composition of fossil
freshwater bivalves (Unioniacea superfamily) of Late Cretaceous–Early
Eocene age collected from six basins in the Laramide tectonic province.
Unionids have relatively thick growth increments aiding the study of
seasonal isotopic variation. Our study examines both the seasonal and
annual δ18O values of ancient river waters as recorded in fossil shells. We
compare these to the δ18O values of modern precipitation and river
water in our discussion of the paleoelevation of ancient river
catchments. We then track the timing and spatial patterns of high
elevation regions in Laramide ranges.
2. Regional setting
Fossil shells were collected from the Alberta foreland basin,
western Williston basin, Crazy Mountains basin, northern Bighorn
basin, Powder River basin and southern Washakie basin (Figs. 1 and 2).
Prior to Laramide deformation, these regions were a broad foreland
basin of the thin-skinned Sevier fold-thrust belt. Laramide deformation partitioned the central Rocky Mountain foreland into discrete local
basins separated by basement-cored uplifts with NE–SW to E–W
orientations (Dickinson et al., 1988). The uplifts are bounded by
moderately dipping to high-angle faults or are broadly anticlinal. The
Crazy Mountains basin, Bighorn basin, Powder River basin and
Washakie basin are of this type. Late Cretaceous through Miocene
age sediments derived from the Sevier thrust belt and Laramide uplifts
were deposited in these basins. Fossil shells from the Washakie basin
are from the Luman Tongue Member of the Green River Formation, a
Fig. 3. Measured and modeled δ18O values of ancient river water for western interior of North America plotted against paleo-latitude of sampling location. Dotted lines are the 1σ
uncertainties of modeled low elevation δ18O values based on paleolatitude data in Table 2 and corrections discussed in text. See text for details of calculations and modeling.
M. Fan, D.L. Dettman / Earth and Planetary Science Letters 286 (2009) 110–121
113
Fig. 4. Seasonal δ18O variation of representative fossil shells in the studied basins. A: Powder River basin; B: western Williston basin; C: Alberta foreland basin; D: Crazy Mountains
basin; E: Bighorn basin; F: southern Washakie basin.
mix of fluvial and lake facies in the early stages of Lake Gosiute (Sklenar
and Anderson, 1985). Shell samples in the other three basins are from
Paleocene to Early Eocene fluvial sediments. In the Alberta foreland
basin, Laramide age deformation is an eastward continuation of thinskinned Sevier fold and thrust, which overthrusted Mesozoic shale and
molasse, but did not generate any basement uplifts (Obsborn et al.,
2006). Shells from the Alberta foreland basin were collected from
sediments of Late Campanian and Maastrichtian age. In the Williston
basin, a depression in the Canadian Shield, fossil shells were collected
from Late Cretaceous and Early Paleocene fluvial sediments. Sample
ages are based on radiometric ages, paleomagnetic ages and intrabasin
mammalian biostratigraphic correlation with magnetostratigraphy
(Table 1).
3. Analytical methods
Aragonite shells used in this study are all unaltered as judged by
physical appearance, cathodoluminescence microscopy and X-ray
diffraction for a subset of the samples. X-ray diffraction was performed
with a Bruker D8 Advance Diffractometer using Cu Kα radiation. When
possible, ten shells and shell fragments from each stratigraphic
horizon were analyzed for stable isotopes. The bulk shell δ18O values
were mostly presented in Dettman and Lohmann (2000), with new
analyses added (supplementary data). Total of 881 individual shell
fossil analyses are included in this paper. Shell Aragonites were
collected by drilling through the shell body, integrating isotopic
variation and yeilding a growth amount-weighted average δ18O value.
The bulk δ18O value presented in this paper refers to the average values
of all analyzed bulk shell samples in stratigraphic horizon except the
Powder River basin, where some data represent single shell analyses.
In addition, we micro-milled 12 selected samples in order to study the
seasonal isotopic variation of the Early Cenozoic rivers. Shells were
sectioned along the axis of maximum growth, mounted as thick
sections (~ 1 mm). Growth bands in cross-sectioned shells were
subsampled using a computer-controlled micro-mill with a 30 μm
sampling resolution (Dettman and Lohmann, 1995). The δ18O and δ13C
values of aragonites were measured using an automated carbonate
preparation device (KIEL-III) coupled to a gas-ratio mass spectrometer
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Seasonal isotopic variation of selected shells from each basin are
used to compare ancient and modern river seasonal variability in the
Laramide region (Figs. 4 and 5). Oxygen isotope ratios of modern river
water are from Coplen and Kendall (2000), modern precipitation are
from the United States Network for Isotopes in Precipitation (USNIP)
(www.uaa.alaska.edu/enri/usnip), Canadian network for Isotopes in
Precipitation (CNIP) (www.science.uwaterloo.ca/~twdedwar/cnip),
and Dutton et al. (2005).
5. Calculation of ancient river water δ18O values
The oxygen isotope composition of Unionid shell aragonite is
controlled by the temperature and δ18O value of the river water in
which it grew (Grossman and Ku, 1986; Dettman et al., 1999). Here, we
use the empirically determined relationship between the bulk δ18O values
of Unionid shell and mean annual river water in temperate climates from
Kohn and Dettman (2007):
18
18
2
δ Oðshell;PDBÞ = ð0:892 0:024Þ δ Oðriver water;SMOWÞ 0:978 0:240ðR = 0:98Þ
ð1Þ
Fig. 5. Seasonal δ18O variation (black circles) of three kinds of river to the east of Rocky
Mountains. A: Milk River; B: Halfmoon Creek; C: Big Sioux River. River water δ18O data
are from Coplen and Kendall (2000). Temperature (gray circles) is the average
temperature for the day of sampling from a nearby climate station (NCDC: http://www.
ncdc.noaa.gov/oa/ncdc.html).
to calculate the mean annual δ18O values of ancient river water. Because
these freshwater bivalves stop growing below approximately 10–12 °C
and their growth is heavily biased to late spring and early summer
temperatures, a large majority of shell aragonite is produced in a limited
range of temperatures (20–25 °C) (Dettman et al., 1999). This growth
temperature bias leads to a good linear correlation between mean annual
river water and bulk shell δ18O values, which can be used to calculate the
mean δ18O value of ancient river water if the climate is temperate and
seasonal (Kohn and Dettman, 2007). If temperatures were not seasonal,
but remained at one temperature extreme (e.g.12° or 30°) throughout the
year, then calculated river water δ18O values could be as much as 2‰ too
high or too low, but this seems extremely unlikely given prominent
growth bands in the shells, indicating seasonal growth cessation, and the
botanical and modeling evidence for moderate seasonality (Wilf, 2000;
Sewall and Sloan, 2006). Our calculated mean annual δ18O values of
ancient river waters range from −23‰ to −5‰SMOW (Fig. 3).
Each shell sample represents an average of several years of growth
in these river systems. Although the average river water δ18O values
calculated from a single shell could be affected by a few years of
(Finnigan MAT 252). Samples (20 to 150 μg) were reacted with
dehydrated phosphoric acid under vacuum at 70 °C. The isotope ratios
are calibrated based on measurements of NBS-19 and NBS-18;
precision is ±0.1‰ for δ18O and ±0.06‰ for δ13C (1σ).
4. Oxygen isotope results
The δ18O values of bulk shells from basins along the front of the
Sevier thrust belt follow a trend from lower values in north (−13.7 ±
1.3‰PDB, Crazy Mountains basin) to higher values in south (−10.7 ±
0.8‰PDB, Bighorn basin, and −7.3 ± 1.5‰PDB, Washakie basin), and the
within-basin variation is small. In the Powder River basin and western
Williston basin, shell δ18O values show large variation between −22.9‰
and −7.9‰ PDB, and −21.8‰ and −9.3‰PDB, respectively. Our results in
the western Williston basin are consistent with values from the eastern
Williston basin (Carpenter et al., 2003; Cochran et al., 2003). The δ18O
values of shells in the Alberta foreland basin vary between −19.3‰ and
−13.3‰ PDB (Fig. 3).
Fig. 6. Regression for sampling station latitude and the δ18O values of river water in lowelevation stations (b200 m) within the USA (data from Coplen and Kendall, 2000).
Open circles are rivers in Hawaii (b23° latitude) or rivers that are sourced in high
elevation catchments with sampling stations at low-elevation, which are not included
in the regression.
M. Fan, D.L. Dettman / Earth and Planetary Science Letters 286 (2009) 110–121
115
Table 2
Paleomagnetic data and paleolatitude.
Sample location
Age
(Ma)
Present latitude
(Nº)
Present longitude
(Wº)
Pole
(Nº)
Pole
(Eº)
A95
P
Paleolatitude
(Nº)
Southward transport
(º)
Alberta foreland basin 1
Alberta foreland basin 2
Williston basin
Crazy Mountains basin
Bighorn basin
Powder River basin
Washakie basin
74–76
68–73
68–62
65–61
54–59
54–58
53–54
49.16
50.95
47.55
46.21
44.84
44.8
40.89
110.41
112.09
107.01
109.73
108.81
106.08
108.66
74.70
74.20
74.20
75.90
75.90
75.90
77.90
204.70
204.80
204.80
196.80
196.80
196.80
179.30
5.90
3.20
3.20
2.90
2.90
2.90
3.40
0.55
0.51
0.59
0.64
0.67
0.68
0.81
58.43
60.75
56.36
53.18
51.67
51.01
43.55
9.3 ± 5.9
9.8 ± 3.2
8.8 ± 3.2
7.0 ± 2.9
6.8 ± 2.9
6.2 ± 2.9
2.7 ± 3.4
Note: A95 is 95% confidence limits. P is the angular distance between sample site and reference pole (Eq.A20 in Butler, 1992). Paleolatitude is calculated from Eq. A26 in Butler (1992).
Southward transport is the latitude difference (and 95% confidence limits) between present latitude and paleolatitude.
anomalous precipitation patterns, the chances of this are reduced by
averaging ten shells from each stratigraphic level.
6. Constraints, corrections, and application to paleoaltimetry
Knowing the δ18O values of ancient river water provides some
insight into the paleoelevation of ancient river catchments, but many
other factors can affect the δ18O of rainfall. Many of these effects (e.g.
temperature, continentality) are combined into the strong relationship
between the δ18O of rainfall and latitude (Dutton et al., 2005). Therefore
we will attempt to remove the latitudinal effect on δ18O values by
comparing the ancient river δ18O values to modern low-elevation rivers
(representing low-elevation rainfall δ18O values) across a range of
latitudes. We use river water δ18O data because it is much more
abundant than precipitation data and it comes closer to a weighted
average of the δ18O of precipitation in the river's catchment. Accounting
for latitudinal differences is particularly important because North
America moved southward, changing latitude by about 10°, during the
time interval under study, and the study region spans another 10° of
latitude. This low-elevation data set will be modified based on studies of
Paleogene meteoric water δ18O values to make the comparison more
appropriate for Late Cretaceous and Cenozoic times. Our data also
provides a check on the latitudinal gradient of δ18O values for low
elevation rainfall in the late Cretaceous and Paleogene.
The δ18O value of modern low elevation (b200 m) river water
across the USA is related to the latitude of river stations (Fig. 6):
18
δ Oðriver water;SMOWÞ = ð0:0012 0:0009Þ j LATj
2
ð0:3940 0:0816Þ jLATj ð2Þ
2
+ 10:2154 1:7811 ðR = 0:91Þ
In the dataset of Coplen and Kendall (2000), the elevations of the
modern river water samples in low elevation (b200 m) are not
necessarily the elevations of river source catchments. Many rivers have
significantly higher source catchments elevation than sampling site, like
Columbia River sampled in Washington, and Missouri River sampled in
South Dakota and North Dakota. Therefore, in our regression we do not
include the rivers that clearly rise at high elevation, and Hawaiian rivers,
which are not applicable to continental interiors. This significantly
improves the regression in Dutton et al. (2005).
The key factors affecting the δ18O values of either modern or ancient
river waters include water vapor source, vapor temperature, source
water δ18O values, latitude, and elevation. The effect of evaporation on
the δ18O value of river water is usually quite small, particularly in humid
climates, and is ignored in this paper. The modern climate of the Rocky
Mountain region and the western Great Plains is influenced by the
competition of three air masses that originate over the Arctic, the Gulf of
Mexico, and the Pacific Ocean (Bryson and Hare, 1974). The Rocky
Mountains, together with the Basin and Range to the west, form a large
high topographic barrier leading to a rain shadow that reduces the
contribution of Pacific-sourced moisture to the Laramide Rocky
Mountain province. To the east of the Rocky Mountains, moisture from
the Gulf of Mexico brings relatively abundant summer rainfall. Prior to
the Latest Cretaceous–Paleocene, the movement of Sevier fold-andthrust belt caused significant shortening and thickening, forming a highelevation hinterland plateau with rugged topographic front to the west
of the Rocky Mountains (DeCelles, 2004; DeCelles and Coogan, 2006),
Given the topographic similarity to today, we assume that large-scale
climate pattern during Early Cenozoic was similar to present and most of
the precipitation is sourced from the Gulf of Mexico region.
Temperature of the sea surface and middle latitude continental
interior in the Late Cretaceous–Early Eocene was higher than present
(Zachos et al., 1994; Wilf, 2000; Zachos et al., 2001). Higher
temperature can influence the δ18O values of both source region
water vapor and water condensates (Dansgaard, 1953), which could
lead to a different stable isotope–latitude relationship than that of
today. One study of the Early Eocene latitudinal-isotopic gradient in
river water shows that the gradient was similar to today in the middle
latitudes, although there is a significant difference in the intercept of
this relationship (Fricke, 2003). The North American Plate has moved
southward at least 5° since the Late Cretaceous (Besse and Courtillot,
2002), which could result in a ~2.4‰ difference between the δ18O
values of precipitation at present and in the Late Cretaceous at a
location (Fig. 6, and Dutton et al., 2005). Neglecting this paleolatitudinal correction will lead to overestimates of the paleoaltitude by
at least 0.8 km (if a lapse rate of 2.8‰/km was applied to this
difference; Poage and Chamberlain, 2001). The paleolatitude of our
sample localities are derived by using the paleopole locations in Besse
and Courtillot (2002), and the dipole equation in Butler (1992). The
lower and upper limits of paleolatitude are derived from the
confidence limits of the paleopole (Α95) (Table 2).
In order to discuss ancient river water δ18O values in terms of stable
isotope patterns seen on the globe today, we use the regression
relationship between the δ18O value of modern low elevation river
water and sampling latitude (Fig. 6), incorporate corrections for 1)
higher intercept due to warmer temperature in the Late Cretaceous–
Early Eocene (+4 ± 2.8‰, Fricke, 2003); 2) changes in the δ18O value of
seawater due to smaller global ice volume (−1‰, Zachos et al., 1994);
and 3) latitude change due to continental drift of North America since
the Late Cretaceous–Early Eocene (Besse and Courtillot, 2002), to model
the low elevation river water 18O values. Errors propagated from the
three corrections yield a δ18O range for low elevation river water (Fig. 3).
If an ancient river catchment was at low elevation, the determined δ18O
values of ancient river water should be in the range of these modeled
δ18O values. Note in Fig. 3 that our most positive δ18O values, probably
sampling the lowest elevation catchments, are within the ancient lowelevation river water field, suggesting that the corrections approximate
low-elevation precipitation relatively well. If the ancient river rises from
a high elevation catchment, the determined δ18O values should be lower
than the area of modeled δ18O values. The elevation of the river source
terrane can then be calculated from the difference in the δ18O value
between the ancient river water and the lower limit of the modeled low
elevation river water, and a lapse rate of −2.8(±0.6)‰/km. This leads to
a conservative estimate of elevation as we use the lower limit of the
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M. Fan, D.L. Dettman / Earth and Planetary Science Letters 286 (2009) 110–121
modeled low elevation river water. This lapse rate is derived from global
precipitation δ18O values (Poage and Chamberlain, 2001), and is similar
to the lapse rate derived from North America precipitation (−2.9‰/km
Dutton et al., 2005). Although our bivalve-based data reconstructs river
water δ18O values, we use an isotopic lapse rate based on precipitation
instead of one based on river water. This is because we will discuss the
elevation of the catchments of these ancient river systems rather than
the elevation of the point in the river system sampled by the bivalves. In
general river sample isotopic lapse rates are much higher than
precipitation-based ones because the former reflect rainfall in the
river catchment at higher elevations than that of the sampling point (e.g.
Dutton et al., 2005). If a river has a high elevation catchment the δ18O
values of the river water will tend to be lower, even if the sample of river
water is collected near sea level. The biggest problem with using the
river water lapse rate is the lack of information on the elevation
difference in a river's catchment and that the elevation of a single
sampling point in a river system has little to do with the overall elevation
of its catchment (Kohn and Dettman, 2007). The shells in this study are
limited to the basin floors, where sediment accumulates and preserves
these fossils. In some basins there is a very large range in the δ18O of river
waters, a range that clearly represents variation in the elevation of the
different river catchments, and not dramatic changes in the elevation of
the basin floor (Dettman and Lohmann, 2000).
The Western Interior Seaway covered a large area of North America
in the Late Cretaceous (Roberts and Kirschbaum, 1995), and this water
body probably dominated the water vapor cycle of the Campanian and
Maastrichtian Alberta foreland and Williston basins. Because this was
basically a coastal system we will not use the low elevation latitude–
rainfall δ18O relationship of Fig. 3, which is more appropriate for a
continental interior region, to interpret the Alberta and Williston
basin data. In this case, we will use a modern empirically-tested model
of the Δδ18O–elevation relationship to reconstruct river catchment
elevations (Rowley et al., 2001; Rowley, 2007). The Δδ18O value is the
δ18O difference between highland precipitation and lowland precipitation. The modeled Δδ18O-elevation relationship, which tracks a
single rising package of water vapor cannot be applied to the Early
Cenozoic samples because of the retreat of the interior seaway and the
possibility that water vapor came into the Laramide region from
different sources. The uncertainties reported in this paper, encompassing the model, corrections, and calculation of paleoelvation all are
one standard deviation (1σ).
7. Seasonal δ18O variation in ancient and modern river waters
The δ18O values of river waters are controlled by surface runoff and
groundwater in river catchments. Rivers with low elevation catchments
tend to have higher δ18O values in contrast to high elevation catchments.
High-elevation snowmelt can feed lowland rivers as surface runoff
during the spring and early summer or year-round as groundwater. The
δ18O values of this type of river water is often very low in spring and early
summer, equivalent to the δ18O values of high-elevation snowmelt
(Horton et al., 1999). At present, in the Laramide region, summer
precipitation amount is about two times that of winter precipitation
amount (Dutton et al., 2005), and the highest river discharge is generally
in March–June (SAGE River Discharge Database, www.sage.wisc.edu/
riverdata). We group the isotopic patterns of rivers in this region into
three types.
1. Rivers originating at high elevation with different sources of water
seasonally. The Milk River flows east from Glacier National Park, Montana, with catchment elevations between 1.0 and 2.8 km. Precipitation
at the closest rain monitoring station (MT-05, Glacial National Park,
968 m) averages −18.3‰SMOW with a range of 15.6‰ (Dutton et al.,
2005; USNIP). The large seasonal δ18O variation in the Milk River at the
Nashua station (48.13°N, 106.36°W, 696 m) shows that the river is fed
by snow melt from elevations higher than station MT−05 in spring
with δ18O values as low as −23‰. In the summer, local and lower
elevation rain seems to dominate, with δ18O value as high as −10‰.
During the winter, when surface runoff is frozen, groundwater supports
the river with intermediate δ18O values (Fig. 5A).
2. Rivers with year-round high-elevation sources of precipitation.
Halfmoon Creek rises in the Sawatch Range, central Colorado,
which contains several of the highest peaks (N4 km) in the Rockies.
At station Malta (39.17°N 106.39°W, 2996 m), the δ18O value is low
and very stable throughout the year, −17.0 ± 1.2‰. There is a small
increase in the δ18O value of river water in the late summer (Fig. 5B).
Precipitation at the closest station (CO-02, Niwot Saddle, 3520 m)
has an annual average δ18O value of −14.3‰ (USNIP). This suggests
that the river is primarily fed by groundwater that integrates annual
precipitation, perhaps derived from higher elevations. Summer local
precipitation seems to increase river water δ18O values only slightly.
3. Rivers originating at low elevation. The Big Sioux River catchment
begins in Roberts County, South Dakota, with an elevation less than
500 m. At station Akron (42.83°N 96.56°W, 341 m), the seasonal δ18O
variation of the river water is very small, −8.6±1.0‰ (Fig. 5C). The
annual average δ18O value of precipitation at the closest station (MN27, Lamberton, 343 m) is −8.6‰ (USNIP). The small seasonal isotopic
cycle and similarity in δ18O value to precipitation suggests that the
river is fed by ground water, which is recharged by local precipitation.
Micromilled shells from the different areas surveyed show a variety
of seasonal oxygen isotopic patterns: three types in the Powder River
and Williston basins, two patterns in the Alberta foreland basin and only
one pattern in the Crazy Mountains, Bighorn and Washakie basins
(Fig. 4). Although the amplitude of seasonal oxygen isotope cycles in
shells are not directly comparable to river water δ18O cycles due to the
seasonal temperature variation, large changes in river water δ18O values
(greater than 4‰ in many cases) must be reflected in the δ18O cycles in
shells. This is because the limited range of shell growth temperatures
can only cause ~4‰ variation in shell δ18O values (Dettman et al., 1999).
In the Powder River basin, shell PR10-86A has a large amplitude δ18O
cycle (−17.9‰ to −7.0‰), similar to that of the Milk River. This highamplitude shell must have lived in a river with large seasonal δ18O
amplitude, suggesting that the river was seasonally fed by snowmelt
from high elevation either as groundwater or as surface runoff. Powder
River basin shell PR2-85 has an isotopic cycle similar to the Big Sioux
River. It has small δ18O amplitude (between −11.8‰ and −7.0‰) and
the mean δ18O value is relatively high. This kind of river is found today at
low elevations, fed by local precipitation. PR3-83 is similar to Halfmoon
Creek, with a small δ18O amplitude, but with a very low mean δ18O value
(−24.3 ± 2.6‰). This kind of river usually originates in high elevations
and is fed by a groundwater supply that is recharged at high elevation
(Fig. 4A). In the western Williston basin, shells with these three types of
seasonal δ18O patterns are present (Fig. 4B). In the Alberta foreland
basin, seasonal δ18O variation in ancient river water recorded in samples
93-13 and Tyrell, is very similar to Big Sioux River, with a small
amplitude and higher δ18O values (−13.6± 1.0‰). In contrast, one
shell, sample 93-11, is similar to Halfmoon Creek, also low amplitude,
but with very low δ18O values (−18.9 ± 0.8‰) (Fig. 4C). In the Crazy
Mountains basin, Bighorn basin, and Washakie basin, the seasonal δ18O
variation in shells have relatively small amplitudes and are intermediate
in average δ18O value. These rivers are of Big Sioux River type, and fed by
local precipitation (Fig. 4D, E, and F). Our Washakie basin shell data are
similar to seasonal records from other well preserved shells from the
Green River basin (Morrill and Koch, 2002).
8. Discussion
8.1. Late Cretaceous High Canadian Rocky Mountains
The δ18O value of paleosol carbonate in the Late Maastrichtian
lower Willow Creek Formation in southern Alberta is −12.1 ± 0.9‰PDB
M. Fan, D.L. Dettman / Earth and Planetary Science Letters 286 (2009) 110–121
117
was active during the Late Cretaceous, and that a high orogenic plateau
sat to the west of the basin (Van der Pluijm et al., 2006).
The δ18O range of river water in the Williston basin is − 18.3 to
−21.5‰SMOW. This is 8.5 to 11.7‰ lower than the δ18O value of the low
elevation river water (Fig. 3) and is similar to that of the Alberta
foreland basin. This suggests mountains with similar elevations, in this
case 4.3 ± 1.0 km, were present (Fig. 7A). Oxygen isotope cycles in
ancient river water further suggest that highland snowmelt fed some
rivers seasonally or yearly (Fig. 4B). Rivers rising in the lowlands of the
basin have much higher δ18O values that fall within the range of
modeled low elevation river water (Fig. 3) and seasonal records of
shell δ18O values reflect both high and low elevation catchments. The
upper Cretaceous and lowest Cenozoic strata in southeastern Montana
contains high amounts of volcanic lithics and andesitic conglomerate,
and paleoflow directions are from the southwest to west in the region;
both suggest that the sediment source terrane was the Sevier thrust
belt in Idaho and southwest Montana (Cherven and Jacob, 1985;
Fastovsky, 1987). Therefore, the high elevation catchments for rivers in
the western Williston basin were most likely in the Rocky Mountains.
Large rivers frequently carry high elevation snowmelt N500 km
without any significant change in the water δ18O value and this is
the most plausible explanation for the low δ18O values in Williston
basin (Fig. 8A). Examples of rivers carrying water with very negative
δ18O values over long distances are the Colorado and Missouri Rivers,
which have significantly lower δ18O values in Arizona and South
Dakota than the local precipitation (Coplen and Kendall, 2000).
8.2. Late Paleocene High surface elevation of eastern Laramide ranges
Fig. 7. A): Estimated paleoelevation of the Late Cretaceous and Early Paloecene
Canadian Rocky Mountains by using thermodynamic based lapse rate in Rowley, 2007;
B): Estimated paleoelevation of mid Paleocene-Early Eocene river source region by
using lapse rate against present latitude of sampling sites in the Laramide Rocky
Mountain basins. Elevation of river catchment with measured δ18O values higher than
modeled low-elevation values are arbitrarily assigned as zero.
(n = 7) (Mack and Cole, 2005). From this we calculate the δ18O value of
lowland mean annual precipitation as −10.3 ± 2.5‰ using the calcite
oxygen isotope fractionation relationship (Kim and O'Neil, 1997) and a
temperature of 12–30 °C, assumed to be the formation temperature of
soil carbonate during warm growing season (Cerling and Quade,1993).
Although evaporation can increase the δ18O values of soil water
relative to mean annual precipitation, this effect is minimal in wet
climates (Quade et al., 2007). Given the proximity of the interior
seaway and wet climate in the Late Cretaceous (Roberts and
Kirschbaum, 1995; Golovneva, 2000), strong evaporation is very
unlikely. The similar δ18O values between the lowland precipitation
and the highest river water suggest that the shells with the relatively
positive oxygen isotope cycle in Fig. 4C lived in river water fed mainly
by local precipitation. The occurrence of much lower values can
therefore be attributed to the low δ18O values of distal highland
precipitation. This can be seen in the shell 93-11 in Fig. 4C, which grew
in water that had low δ18O values all year long. Numerous other shells
with bulk δ18O values in the −18‰ to − 20‰ range have similar cycles
and these low values cannot be the result of seasonal low-elevation
snowmelt events. Our estimated highland elevation is 4.3 ± 1.0 km
during Late Cretaceous (Fig. 7A). In the Alberta foreland basin,
paleocurrent directions are generally east to southeast (Mack and
Cole, 2005), suggesting that high elevation snowmelt came from the
Canadian Rocky Mountains to the west of the basin. This is consistent
with tectonic reconstructions of this region, which argue that thrusting
In the Powder River basin, the lowest δ18O value of the river water is
−22‰SMOW, which is ~ 11‰ below the minimum δ18O value modeled
for low-elevation ancient river water (Fig. 3). The negative offset
reflects high elevation in the river catchments. Seasonal oxygen
isotope variation in ancient shells further suggests the presence of both
highland and low elevation precipitation in local rivers (Fig. 4A). The
estimated elevation of the mountain ranges that provide runoff to the
Powder River basin is 4.5 ± 1.3 km in the Late Paleocene (Fig. 7B). The
simplest scenario for the paleography in the Powder River basin based
on these three different types of river isotopic records would be a
variety of river catchments that carried distal high elevation waters
and/or basinal lowland precipitation (Fig. 8B).
The high elevation region in the eastern Laramide province was
most likely the Bighorn Mountains in northern Wyoming and Black
Hills in South Dakota. Today the rivers of this region flow east and
northeast from the high Laramide ranges in northern Wyoming and
southwest Montana. These rivers have low δ18O values due to high
elevation snowmelt input. For example, the Tongue River which flows
northward from the northern Bighorn Mountains has a δ18O value of
−16.6 ± 2.0‰SMOW; tributaries of the Powder River with δ18O values
of ~−17‰ flow from the eastern Bighorn Mountains and western Black
Hills toward the center of the basin merging as the Powder River
(Coplen and Kendall, 2000). The Powder River and Tongue River then
join the Yellowstone River in southwestern Montana, and form a large
tributary of Missouri River. A Late Paleocene scenario similar to the
modern drainage pattern could explain the very low δ18O values of the
river waters in the Powder River basin. Such a paleodrainage system is
supported by the paleogeographical reconstruction of the Powder
River basin, which is very similar to the pattern of the modern drainage
system just described (Flores and Ethridge, 1985). Low-temperature
thermochronology studies also support this paleogeography, suggesting major exhumation of the Bighorn Mountains at 65 ± 5 Ma and
Early Paleocene cooling of the Black Hills (Strecker,1996; Crowley et al.,
2002). In addition, Late Paleocene–Early Eocene synorogenic conglomerates with a high proportion of Precambrian basement clasts are
found along the eastern flank of the Bighorn Mountains (Hoy and
Ridgeway, 1997).
118
M. Fan, D.L. Dettman / Earth and Planetary Science Letters 286 (2009) 110–121
Fig. 8. Paleodrainage reconstruction of the studied area based on the presented data, and the inferred elevation of the Rocky Mountains. Lines with arrows stand for the paleorivers
with the δ18O value shown. See Fig. 1 for names of major structures and basins, and references of paleoflow directions.
Fricke (2003) pointed out that seasonal or episodic Arctic air mass
incursions could result in very low δ18O values in the Powder River
basin. Although this phenomenon can explain very negative δ18O
values in seasonal precipitation (snow) at low elevations today, it is
very unlikely to affect the average annual or warm season δ18O value
of river waters. In southern Saskatchewan, where the present latitude
is equivalent to the paleolatitude of the Williston basin and Powder
River basin, winter snow δ18O values are often less than −25‰
(CNIP), but the modeled δ18O value of low-elevation river water is
−16 ± 2‰, which is the same as the δ18O value of the local mean
annual precipitation (CNIP). In this case the low-elevation rivers are
not biased in favor of winter precipitation, rather they are supported
by local groundwater. This argument is also supported by the fact that
low elevation rivers in Minnesota and Michigan have δ18O values
similar to that of the local mean annual precipitation, which is
relatively high (−8 to −10‰), even though these areas are equally
subject to significant water input from arctic air masses (Coplen and
Kendall, 2000; USNIP).
8.3. Lower surface elevation of the western Laramide province before the
Early Eocene
The δ18O values of Paleocene and Early Eocene river water in the
Crazy Mountains basin, Bighorn basin, and Washakie basin fall within or
close to the range of modeled low elevation river water (Fig. 3), which
suggests that the catchments of these ancient rivers were at relatively
low elevation. Seasonal oxygen isotopic variation in shells from the three
basins further shows that river water did not have extremely low δ18O
values (b−15‰) seasonally or annually (Fig. 4D, E, and F). Although we
suggest that this indicates low elevation catchments for the sampled
rivers in these localities, we can not completely rule out some effect of
local relief on these samples for two reasons: 1) both modeled and
measured low elevation river water δ18O values have a range of ~4.5‰,
which could mask ~2 km of elevation variability; 2) the δ18O values of
ancient mean annual precipitation recorded in the paleosol carbonate
nodules of Paleocene–Early Eocene age in this area are higher than our
calculated river water δ18O values, which suggests an elevation
difference between local and distal precipitation sources. The δ18O
values of paleosol carbonates in the Wind River basin are −10.0‰ to
−8.0‰PDB in the Early Eocene (our unpublished data), and are −10.0‰
to −7.5‰PDB in the Bighorn basin in the Late Paleocene–Early Eocene
(Koch et al., 1995). Using a broad temperature range (12–30 °C) we can
estimate the δ18O values of rainfall in these basins as between −4.5 and
−11‰SMOW. Unionid shells in Bighorn basin (−12.6‰ to −8.6‰PDB)
imply that water δ18O values are in the −8.5 to −13‰SMOW range
(using Eq. (1), above). In the Sage Creek basin, southwest Montana, the
δ18O values of paleosol carbonates varies between − 9.0‰ and
−5.0‰PDB during Early Eocene Climatic Optimum, while the δ18O
values of Paleocene fluvial calcite cements are −11.0‰ to −9.0‰PDB
(Kent-Corson et al., 2006), again this suggests that river water δ18O
values are lower than local precipitation. This difference is increased if
the Eocene soils were warmer than the Paleocene rivers. The differences
between the δ18O values of ancient river carbonates and soil carbonate
in these basins suggest a maximum local relief of 1–2 km. Such relief is
tectonically very plausible because crustal shortening and thickening of
the Sevier fold-thrust belt during the Late Cretaceous–Paleocene formed
a high-elevation hinterland plateau to the west of the Rocky Mountains
(DeCelles, 2004; DeCelles and Coogan, 2006). Paleocene paleocurrent
directions in the studied strata of the Bighorn and Crazy Mountains
basins are generally eastward, which supports the scenario that rivers
sourced in the Sevier thrust belt delivered water with lower δ18O values
to the basins (Dickinson et al., 1988; DeCelles et al., 1991; Borrell and
Hendrix, 2000). There are no robust paleocurrent direction measurements in the Luman Tongue Member in the Washakie basin. Our data
therefore suggest there was no Laramide range in the western Laramide
province as high as those in the eastern region in the Late Paleocene.
M. Fan, D.L. Dettman / Earth and Planetary Science Letters 286 (2009) 110–121
119
Fig. 9. Schematic cross-sections showing the mechanism of forming high ranges in the eastern Laramide province earlier than the western province. Not to scale.
8.4. Implications
This study adds to a growing body of knowledge regarding the
topographic development of the Rocky Mountains. High elevation, 4.5 ±
1.3 km, developed in the eastern Laramide province, and a catchment
draining the high regions of the Bighorn Mountains and the Black Hills
formed in the Powder River basin by the Late Paleocene. Our data from
the western Laramide province (Crazy Mountains, Bighorn, and
Washakie basins) suggest that a maximum of 1–2 km of local relief
existed between the Sevier thrust belt or Laramide ranges and basins in
the Paleocene–Early Eocene. Therefore, ranges in eastern Laramide
province reached high surface elevation earlier than ones in west. This
pattern is analogous to the westward migration of the exhumation front
in the Sierras Pampeanas and the high surface elevation of the Sierra
Grande in the eastern Sierras Pampeanas (Coughlin et al., 1998).
Significant elevation increase of Laramide ranges after the Late
Paleocene in the western Laramide province is required to form the
present landscape of the Rocky Mountains.
Sedimentological and structural evidence have shown that Laramide
deformation was initiated in the Late Cretaceous (e.g., Gries, 1983 and
ref. therein, Dickinson et al., 1988; Bird, 1998; DeCelles, 2004). However,
this study shows that Laramide ranges with elevations comparable to
today's were not formed until the Late Paleocene in the eastern Laramide
province, and that ranges in eastern province attained high topography
prior to the western province. Erslev (1993) suggested that the Laramide
uplifts are a system of thrusts and back-thrusts which were connected to
a detachment in the lower curst. We therefore propose that the pattern
we observe could be explained by southwestward progression of backthrusts soled into an east-directed master detachment that formed
earlier (Fig. 9) (Erslev, 1993, 2005). The shallow subduction of the
Farallon Plate beneath North America buckled the lithosphere in the
Sevier foreland by basal traction or horizontal endloading of highlands
in the Sevier hinterland during the Late Cretaceous to mid Paleocene
time (Bird, 1998; Tikoff and Maxson, 2001). This buckling could have led
to high Laramide ranges relative to the basins. The subsequent
development of an east-directed crustal detachment may have formed
the Black Hills and Bighorn Mountain uplifts in the Late Paleocene at the
detachment tips. Following this, back-thrusts could develop in the
folded area and lead to high ranges in the western Laramide province
(Erslev, 1993, 2005). The development of the detachment thrust and
back-thrusts may follow the magmatic sweep caused by eastward
shallow subduction and westward rollback of the subducted slab
(Constenius, 1996).
Although our research is based on what is by far the largest paleosurface water δ18O data set collected to date, with good age control
and multiple basin distribution, it is clear that a much more detailed
temporal and spatial data set is needed to reconstruct the elevation
history of the Laramide ranges and test our proposed kinematic
model. In addition, more work is needed on the post-Laramide basins
to document later erosional exhumation, uplift, or subsidence that led
to the topographic patterns we see today.
9. Conclusion
After removing the latitudinal δ18O effects on the δ18O of ancient river
water, we can discuss the patterns of elevation change across the Late
Cretaceous–Early Eocene foreland basin of the Canadian Rocky Mountains
and the intermontane basins of the central Laramide Rocky Mountains.
The Canadian Rocky Mountains were 4.3±1.0 km high during the Late
Cretaceous, and snowmelt from the Rocky Mountains recharged the river
water in the Alberta foreland basin, and western Williston basin both
annually and seasonally. Laramide uplifts attained elevations up to 4.5±
1.3 km in northeast Wyoming in the Late Paleocene, and the most likely
locus of this high elevation was the ancestral Bighorn Mountains and the
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M. Fan, D.L. Dettman / Earth and Planetary Science Letters 286 (2009) 110–121
Black Hills. Paleodrainage patterns similar to the present developed in the
Powder River basin in the Late Paleocene; rivers in the basin had
catchments at a variety of elevations, some carrying high elevation
precipitation, others sourced at low elevations. No comparably high
topography developed in the western Laramide province before the Early
Eocene. This east–west difference in timing of high elevation development for the Laramide ranges can be explained by southwestward
progression of back-thrusts soled into an earlier northeast-directed
master detachment, which may be formed by the westward rollback of
a subducted shallow slab.
Acknowledgements
The bivalve fossil collections in this paper are largely based on the
field notes of the late John H. Hanley. We would like to thank Jay Quade,
Peter DeCelles and Paul Kapp for their helpful discussion. This manuscript was significantly improved by reviews by Carmala Garzione, and
an anonymous reviewer.
Appendix A. Supplementary data
Supplementary data associated with this article can be found, in
the online version, at doi:10.1016/j.epsl.2009.06.024.
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